Four sites (1097, 1100, 1102, and 1103) were drilled on the outer continental shelf during ODP Leg 178, and three sites (1095, 1096, and 1101) were drilled on sediment drifts on the upper continental rise (located in Fig. F2 and described in Barker, Camerlenghi, Acton, et al., 1999). The continental shelf and rise are the two main depositional environments identified by ANTOSTRAT as containing an accessible record of long-term glacial history. In addition, two sites (1098 and 1099) were drilled within Palmer Deep, an inner-shelf basin south of Anvers Island, where an ultra high resolution postglacial (Holocene) sedimentary record is preserved. This paper does not address the results of Palmer Deep drilling, which is the subject of a separate synthesis (Domack, Chap. 34, this volume).
The main aims of Leg 178 drilling, sampling shelf and rise environments for information on glacial history, confirming models of margin evolution, validating the drilling strategy, and identifying some of the technical problems involved, were successfully achieved. It was shown that a glacial environment existed along the Antarctic Peninsula throughout the last 9-10 m.y., with glacial deposition on both shelf and rise but a change at ~4.5 Ma to progradation of the outer continental shelf (Barker, Camerlenghi, Acton, et al., 1999). Predictably, the continuously deposited and more readily recovered sediments of the continental rise have attracted much more attention postcruise than those of the continental shelf.
The depositional parts of the high-latitude (glacial) outer continental shelf and slope are underlain by a prograded wedge, most probably largely composed of diamict. This was mostly transported as a low-viscosity, sheared, overpressured basal till beneath low-profile ice streams (the major means of ice sheet drainage) that were grounded to the shelf edge during glacial maxima and was deposited in "trough-mouth fans" at the outlets of major ice drainage areas as topsets on the outer part of the (generally overdeepened and inward sloping) shelf and as foresets on the uppermost slope (Alley et al., 1989; Larter and Barker, 1989, 1991b; Vorren et al., 1989; Boulton, 1990; Bartek et al., 1991; Cooper et al., 1991; Pope and Anderson, 1992; Pudsey et al., 1994; Vanneste and Larter, 1995). During interglacials, the ice sheet grounding line is well inshore and deposition on shelf and slope is slow, comprising mainly fine-grained hemipelagic and diatomaceous muds and oozes. Subsequent ice streams may erode previously deposited shelf sediment, so the shelf record is typically incomplete. On the Antarctic Peninsula Pacific margin between ~63° and 68°S, Pliocene-Pleistocene deposits (the progradational Unit S2 and overlying, usually less progradational Unit S1 of Larter and Barker, 1989) are focused into four lobes (Larter and Cunningham, 1993), within which both topsets and prograding foresets are thickest; Unit (more precisely, sequence group, since S1, S2, and S3 each equate to deposition during many glacial-interglacial cycles) S1 may not be continuous between all lobes. The Antarctic Peninsula continental slope is unusually steep in both the lobe and the interlobe areas (Fig. F2). While most workers assume that deposits within the lobes were transported there by ice streams, which may diverge or migrate in order to allow deposition to take place over a wider area (the conventional view of glacial deposition on shelf and slope), Rebesco et al. (1998) now consider, on the basis of shelf bathymetry, that ice streams flowed across the interlobe areas, transporting the bulk of glacial sediment across the shelf with relatively little deposition there or on the slope, and that the lobes were formed by sediment transported on and deposited from slower-flowing ice-stream flanks. This divergence of view persists but is not addressed further here since it has little or no effect upon the value of shelf, slope, and rise sediments as recorders of glacial history.
Three Leg 178 sites (1100, 1102, and 1103) lie on a transect along the axis of Lobe 1 (Fig. F2), and one (Site 1097) was located between Lobes 3 and 4, where access to deeper layers was better (Shipboard Scientific Party, 1999a, 1999c). All sites were drilled using the rotary core barrel (RCB). Penetration was low at Sites 1100 and 1102. At the other sites, recovery was very low within the upper layers, where the finer-grained matrix of the diamicts was insufficiently consolidated to support the clasts during drilling. Even at greater depths, where the diamicts were more consolidated, recovery improved only to a maximum of 34% in the lower 115 m of Site 1103.
The geometry of the glacial shelf deposits and the achievements of shelf drilling may be summarized with respect to Figure F3, a seismic reflection profile along the axis of progradational Lobe 1, the line of the Shelf Transect (Shipboard Scientific Party, 1999a), with the effective penetration of Site 1097 (from an interlobe area) projected onto it. Briefly, penetration at Site 1100 was only 110 meters below seafloor (mbsf), entirely within sequence group S1 topsets, and recovery was very low. At Site 1102 on the continental shelf edge, penetration was even less (only 15 mbsf), but a camera survey (see the video recording in Barker, Camerlenghi, Acton, et al., 1999), undertaken during a swell-induced pause in drilling, revealed a boulder carapace on the seabed that we interpreted as the result of bottom-current winnowing of fine-grained sediment following disturbance of diamicts by iceberg keels. This concentration of coarse material, we thought, while inhibiting penetration and recovery, might explain the strong seismic reflectors seen at many paleoshelf breaks.
Figure F3 shows that drilling at Site 1103 penetrated seismic sequence group S3 (to 363 mbsf) below an unconformity at the base of sequence group S1; sequence group S2 is absent. In contrast, sequence group S3 was sampled at Site 1097 (to 437 mbsf) below an apparently conformable boundary with sequence group S2 and thin sequence groups S1 and S2. Before drilling, sequence group S3 was considered "preglacial" by some (e.g., Larter et al., 1997), but drilling showed its glacial nature at both sites. Its base was not reached at either site.
Sequence group S3 has the appearance that it shows in Figure F3 along the entire margin. It shows parallel to gently divergent bedding on seismic reflection profiles, and preliminary indications are that, unlike the overlying sequence groups S2 and S1, it is not focused into lobes. Dip of sequence group S3 is horizontal in the vicinity of Site 1097 and gently seaward around Site 1103 (Fig. F3). If the sequence group S3 bedding is considered as shelf topsets (and the bedding dip in Fig. F3 results from postdepositional tilt), then the sequence group lacks foresets and is aggradational. In this it contrasts markedly with the overlying, clearly progradational sequence group S2. On many profiles along the margin it is possible to identify a transition to foreset development, with topset truncation in places, formally placed at the base of sequence group S2, as in Figure F3. On board ship (Shipboard Scientific Party, 1999a, 1999c), the state of preservation of benthic foraminifers and other shell material was used to distinguish depositional environments for sequence group S3, ranging from subglacial through proximal proglacial to glacial marine at Site 1097, taken as indicating a continental shelf environment. At Site 1103, the same methods showed no evidence of subglacial deposition (a glacial shelf indicator) and the sequence group S3 sediments showed evidence of sorting and downslope movement, leading to a division of opinion on board ship between a shelf and an upper-slope paleoenvironment.
Throughout the continental shelf holes, even within sequence groups S2 and S1, a sparse admixture of biogenic material, mainly diatoms, allowed some age attribution despite poor recovery and some reworking. It seemed likely that, on the Antarctic Peninsula shelf at least, the ice-base processes responsible for terrigenous sediment transport during glacials had usually involved also a degree of erosion and admixture of the biogenic and fine-grained hemipelagic sediment deposited during preceding interglacials. If the microfossils were relatively fresh and whole, making identification easier, then both the time interval between original deposition, erosion, and redeposition and the transport path seemed likely to have been short. In the context of the relatively high-energy environment of the glacial continental shelf, such dating can be considered precise. Thus, the sequence group S3-S2 transition at Site 1097 (considered conformable on the evidence of seismic reflection profiles), was dated as lying within the Thalassiosira inura diatom zone, at ~4.5 Ma (although if the sequence group S3/S2 boundary is located beneath the deepest evident foreset, this transition could be slightly older). The lower drilled part of sequence group S3 at Site 1097 and all of sequence group S3 sampled below the S3/S1 unconformity at Site 1103 was considered to lie within the Actinocyclus ingens v. ovalis diatom zone (6.3-8.6 Ma). It seems reasonable to conclude that S3 is a synchronous depositional sequence group, of which the uppermost part was sampled at Site 1097 but not at Site 1103, as Figure F3 suggests. The uniform, parallel-bedded character of sequence group S3 on seismic reflection profiles argues against it representing two contrasting depositional environments. Eyles et al. (2001) propose that the sorting detected in some sediments at Site 1103 reflects a continental slope environment, but many others have recorded gravity sliding (resulting in some sorting) on the glacial shelf, down the nose of a "till tongue" ("till delta" or "subglacial delta"), for example (Alley et al., 1989; Vorren et al., 1989; Bart and Anderson, 1995; Vanneste and Larter, 1995). These considerations and the partly subglacial environment of deposition of sequence group S3 sediments at Site 1097 suggest to us that sequence group S3 was everywhere a shelf deposit. It was concluded also that sequence group S3 reflected in some way a less glacial climate than that under which the succeeding sequence groups S2 and S1 were deposited, but it is difficult to be confident about this, given the extremely poor recovery at shallow depths; certainly the recovered sequence group S3 sediments showed a wider range of environments, some of them in a sense less glacial than those of recovered sequence group S2 and S1 sediments, but this is capable of more than one explanation.
The very poor recovery also ensured that the sequence group S2/S1 boundary was not precisely dated at Site 1097. The available diatom ages from sequence group S1 were Quaternary and possibly late Pliocene and from sequence group S2 were Pliocene (Barker, Camerlenghi, Acton, et al., 1999) (Fig. F4).
Postcruise work on data and samples from the drilling leg has been devoted to improving the age constraints on sequence group S3 sediments and the precision with which data from the drill sites could be used to constrain the ages and natures of seismic sequence groups identified in the very large seismic reflection data set from the shelf. In addition, Camerlenghi et al. (in press) undertook a decompaction and backstripping experiment on the sediments along the line of the shelf transect (Fig. F3), the axis of deposition of progradational Lobe 1. This experiment, considered preliminary by the authors, is a useful attempt to assess quantitatively the flexural response to sediment loading, thermal subsidence from ridge-crest subduction, and sediment compaction under non-hydrostatic load. The authors conclude that the present overdeepened and inward-sloping shelf, considered to indicate unequivocally a fully glacial regime, developed only within the time of deposition of sequence group S2 and suggests therefore that the environment of deposition of the older sequence group S3 was indeed less glacial. They find these conclusions robust to a range of flexural, thermal, and compaction parameters. The analysis has other implications: neither the parallel bedding of sequence group S3 nor the topsets of the presumed topset/foreset couples of the lower part of sequence group S2 (see Fig. F3) are raised to the horizontal within the model; little of their present-day dip is eliminated. This would argue against interpretation of the depositional environment of sequence group S3 as continental shelf. Moreover, also it rules out the topset/foreset interpretation of the basal sequence group S2 features, which are common along the margin and in all other respects resemble features that are characteristic of the sequence group S2/S1 boundary and are common within sequence group S1. One possible explanation lies in the model's assumption of uniform, noncompactible material beneath sequence group S3 all the way along the model profile. To the southeast (inshore) this material is the high-velocity mid-shelf high, where the assumption almost certainly holds, but to the northwest (offshore) it becomes the precollision sequence S4 sediments and their more distal sedimentary equivalents of the precollision accretionary prism. Many accretionary prisms have low velocity and low density toward the trench (e.g., Cochrane et al., 1996) and would compact significantly under load. We should perhaps await more detailed modeling before drawing firm conclusions about the depositional environment of sequence groups S2 and S3 and the implications for paleoclimate.
The rather sparse shipboard constraints on the age of sequence group S3 have been augmented postcruise by two different investigations and by additional work on microfossils (M. Iwai and L. Osterman, pers. comm., 2001) (Fig. F4). The assignment of the depth range 320-355 mbsf at Site 1103 to the diatom A. ingens v. ovalis Zone (6.3-8.6 Ma) is supported in general terms by Sr isotopic ages of 7.4 and 7.8 Ma on barnacle fragments from 262 to 263 mbsf (Lavelle et al., Chap. 27, this volume) and by the youngest (7.6 Ma) of a range of 40Ar/39Ar ages for volcanic glass fragments from 337 mbsf (Di Vincenzo et al., Chap. 22, this volume). Both ages are subject to uncertainties; the barnacles are well preserved and their ages cluster well, but preservation of the fragments is slightly poorer than that of a single fragment from one of the samples that gave an older age. The argon age is merely the youngest of a wide range of ages, each from a group of several small grains of volcanic glass present in low concentrations within the sediments and showing signs of mechanical reworking, and is unsupported by studies on potassium and argon mobility. Nevertheless, the close coincidence of all these ages supports the view that much of the sampled part of sequence group S3 is of late Miocene age. A third study of pollen and spores and nannofossils from the shelf sites (Iwai et al., Chap. 28, this volume), has yielded mainly rare forms, unfortunately probably recycled and of little stratigraphic or paleoclimatic value.
Two studies were aimed at making more precise use of data from the drill sites for calibration of the large seismic reflection data set from the continental shelf. At Site 1103, the seismic velocity information provided by downhole logging was of doubtful quality initially, but Moerz et al. (Chap. 19, this volume) have undertaken a painstaking reanalysis of the data in a successful attempt to eliminate instrumental uncertainties. Tinivella et al. (Chap. 16, this volume) describe tomographic studies of seismic velocity at five points along seismic reflection profile I95- 152, which coincides with the shelf transect (Fig. F3). The tomographic analysis shows a large velocity increase with depth within topsets and high velocities in underlying foresets. In general, the results suggest that velocity is more strongly related to burial depth than to age or depositional environment, but there are exceptions; in one place (not drilled but equivalent to the section drilled at Site 1097) they show a low velocity within the upper part of sequence group S3, and, at the very shelf edge, foresets have much higher velocities than topsets, even at shallow depth. Where the study can be compared with the downhole log analysis from Site 1103, it shows a strong velocity gradient within the topsets that is not seen in the log data. The optimal correlation of drilling results with seismic reflection profiles is probably made by combining the Site 1103 downhole log velocity data down to ~233 mbsf (below which the only velocity measurements are on sparse shipboard samples) with the tomographic velocities beneath. A study of the large seismic reflection data set from the continental shelf, using the results of drilling, is under way but could not be completed for this volume.
A feature of high-latitude glacial sedimentation is that the fine-grained part of the unstable component of sediments deposited from grounded ice on the uppermost continental slope can accumulate within sediment drifts on the upper continental rise (Kuvaas and Leitchenkov, 1992; Tomlinson et al., 1992; McGinnis and Hayes, 1995; Rebesco et al., 1996, 1997; Clausen, 1998), following slumping and turbidity current flow down the slope and suspension as a nepheloid layer within bottom currents. Provided that bottom currents are slow and that the slope residence time of this unstable component of uppermost slope deposits is short compared with a glacial cycle (that is, the instability involves only small-scale mass wasting), the drift sediments can provide an indirect but continuous high-resolution and recoverable record of glacial history. A lesser interglacial sea-ice cover permits primary biogenic production, hence pelagic deposition and dating. Off the Pacific margin of the Antarctic Peninsula these conditions appear to hold; current meter measurements (Camerlenghi et al., 1997a) have revealed relatively slow bottom currents (contour-parallel but generally southwest flowing with mean speed of 6.2 cm/s and not exceeding 20 cm/s over a 10-month period in 1995), there is indirect evidence (e.g., Pudsey, 1992) that glacial-age currents were no faster, and the character of the upper continental slope (slump headwalls and gullying seen on deep-tow boomer profiles—Vanneste and Larter, 1995) indicates that recent instability has been small scale. Both piston cores (Camerlenghi et al., 1997b; Pudsey and Camerlenghi, 1998; Pudsey, 2000; Lucchi et al., in press) and the preliminary results of Leg 178 drilling have verified the preserved and dateable record in the drifts of late Pleistocene glacial-interglacial variability.
Most of the drifts (including those drilled) are separated from the continental slope and from each other (Fig. F2) by a dendritic pattern of channels originating at the base of slope (e.g., Tomlinson et al., 1992), with axes up to 1 km deeper than the drift crests. The channels are maintained by turbidity currents, which carry the coarse silt- and sand-sized components of the unstable upper-slope sediments to the abyssal plain. These drifts are considered an end-member of the drift category, since their topographic distinction from the channels is maintained (it is thought) not by bottom currents (which here are weak), but by erosion by the turbidity currents themselves. Thus, it is possible for interglacial, mainly pelagic sediments to be eroded and resuspended by the turbidity currents on the rise itself (within the channels) and incorporated into the drifts. Rebesco et al. (in press) recently summarized all of the available marine geophysical and marine geological data from the Antarctic Peninsula sediment drifts. Of particular interest is the swath bathymetry in the northeastern area (Drifts 1-3 and part of Drift 4) mentioned originally by Canals et al. (1998), which supersedes the GLORIA side-scan data reported by Tomlinson et al. (1992) and clearly shows both the essentially asymmetric drifts and intervening channels.
Three sites on the continental rise were drilled during Leg 178 (Fig. F2) (Barker, Camerlenghi, Acton, et al., 1999). Two sites together examined a single drift (Drift 7 of Rebesco et al., 1996). Site 1096 was the closer to the continental shelf, in 3152 m water depth, and sampled an expanded section extending back to 4.7 Ma by combined use of the advanced hydraulic piston corer (APC) and extended core barrel (XCB) to 608 mbsf, with repeated APC coring at shallow depth in order to obtain a more complete section. Site 1095, in 3842 m water depth at a more distal location, where the lower section was more accessible because the upper section was thinner, sampled to ~10 Ma at a maximum depth of 570 mbsf, again with repeated APC coring to improve recovery in the shallow section. At both sites the deep hole was logged almost to maximum depth. This dual-site strategy avoided drilling through a diffuse bottom-simulating reflector seen on seismic reflection profiles at ~700 ms depth and interpreted as caused by silica diagenesis (Barker and Camerlenghi, 1999), thus perhaps obliterating siliceous biostratigraphy. The third site (Site 1101), in 3280 m water depth, sampled a different drift (Drift 4 of Rebesco et al., 1996) back to 3.1 Ma in a single APC/XCB hole to 218 mbsf.
Recovery at Sites 1095 and 1096 was high, after taking into account the multiple coring of the shallow section: Site 1096 recovery = 88%, Site 1095 recovery = 94% down to ~484 m (where recovery dropped sharply), or 82% overall. Recovery at Site 1101 was even higher, 99%. It was possible to obtain a magnetostratigraphic record on board ship from cores and logs and to examine the preserved microfossils. Postcruise magnetostratigraphic studies have been undertaken in more detail, both on U-channel samples of the core (Acton et al., Chap. 37, this volume) and by means of a rigorous reexamination of the Geological High-Resolution Magnetic Tool (GHMT) log data (Williams et al., Chap. 31, this volume). In addition, Brachfeld et al. (Chap. 14, this volume) have examined the magnetic mineral assemblage as a contribution toward high-resolution magnetostratigraphic studies such as the detailed U-channel study of part of the Matuyama Chron at Site 1101 (Guyodo et al., 2001).
The biostratigraphic record was more difficult to interpret because of the likelihood that microfossils had been reworked from the shelf and rise before deposition within the drifts, as a result of the dominance of processes of reworking on the continental shelf during glacial periods and gravity sliding with resuspension on the continental slope. Shipboard efforts have been supplemented by additional studies of the main microfossil groups (pollen and spores and nannofossils by Iwai et al., Chap. 28; radiolarians by Lazarus, Chap. 13; dinoflagellates by Pudsey and Harland, Chap. 2; diatoms by Winter and Iwai, Chap. 29 [supplemented by a report of identification and occurrence at several sites— Iwai and Winter, Chap. 35]; and nannofossils by Winter and Wise, Chap. 26, all this volume). The shipboard work, many of these studies, and other work have been brought together into a magnetobiostratigraphic synthesis that focuses mainly on Site 1095, entirely on the rise drifts (Iwai et al., Chap. 36, this volume). Not all inconsistencies have been fully resolved in this work, but it is likely to stand as a stratigraphic consensus for some time, affecting all future and unfinished work on samples and data from the drift sites. The revised correlation between multiply cored intervals and the calculation of meters composite depth (mcd) at Sites 1095 and 1096 (Barker, Chap. 6, this volume) is similar in retaining uncertainties but is likely to come into common use.
In the magnetobiostratigraphic synthesis (Iwai et al., Chap. 36, this volume), the magnetostratigraphy of the continental rise sites (1095, 1096, and 1101) has been revised as a result of U-channel measurements (described in detail by Acton et al., Chap. 37, this volume) and (for Sites 1095 and 1096) the spliced mcd scale at shallow depths, and merged with the combined results of the wide range of shipboard and postcruise biostratigraphic studies. Several key discrepancies reported within the Leg 178 Initial Reports volume have been reexamined, and either a resolution or a most likely path to resolution has been agreed upon. Two developments are particularly significant. First, drilling at Site 1095 was considered to have encountered a hiatus in sedimentation at ~60 mbsf (Shipboard Scientific Party, 1999b) on the basis of the seismic reflection profile through the site and missing magnetic reversals. However, the revised magnetic reversal stratigraphy is much more complete, and the relatively low resolution of the seismic reflection profile is able to accommodate either continuous deposition or a brief hiatus (of 100-200 k.y., for example, as now proposed). Second, there was within the shipboard stratigraphy described in the Leg 178 Initial Reports volume an unresolved discrepancy between diatom (also, subsequently, radiolarian) stratigraphy and the magnetic reversal record at around the Miocene/Pliocene boundary at Site 1095. Following reexamination of all data sets, the magnetic reversal stratigraphy is accepted and the discrepancy is now considered to require most probably a reexamination of published Southern Ocean biosiliceous zonation with particular focus on the occurrence of depositional hiatuses at many sites. The revised age-depth curve for Site 1095 is shown as Figure F5.
In addition to the stratigraphy, Iwai et al. (Chap. 36, this volume) also confirm the presence of a diverse assemblage of neritic and benthic diatoms, with particular abundance of the neritic diatom Paralia sulcata in upper Miocene (7.5 to 6.7 Ma and slightly younger) sediments at Site 1095, being taken to indicate the existence at that time of a wide and shallow (<100 m) continental shelf. This is within the time of deposition of sequence group S3 sediments on the shelf. P. sulcata is also abundant in recovered sediments deposited at Site 1097 on the continental shelf over the same period (M. Iwai, pers. comm., 2001). These occurrences support the conclusion of the decompaction model described above (Camerlenghi et al., in press), that initial glacial overdeepening took place only after the start of deposition of sequence group S2.
Shipboard sedimentologic studies identified a cyclicity in deposition on the drifts that for the late Quaternary could be associated with glacial cycles (Fig. F6); during glacials, with the grounded ice sheet at the continental shelf edge and (presumably) persistent sea-ice cover, drift sedimentation was relatively rapid and comprised gray, terrigenous, finely laminated barren silty clays. Interglacial sediments were thin, brown, diatomaceous, and bioturbated silty clays. Ice-rafted detritus (IRD) was more evident in the bioturbated clays, but the differences in sedimentation rate through the glacial cycle hindered an assessment of absolute rates of ice rafting. At the shallower sites, nannofossils and foraminifers were preserved, at least in the Stage 5 interglacial. This character accords with that seen in piston cores from the drifts (e.g., Pudsey and Camerlenghi, 1998; Pudsey, 2000). Below ~30-50 mbsf at the sites (0.8-2.0 Ma), the alternations of barren, laminated with massive bioturbated and microfossil-bearing beds persist but are less regular. At the distal site (1095) in particular, there are abundant thin silt and mud turbidites within the glacial sequences, and at the two shallower sites (1096 and 1101), nannofossils and foraminifers are preserved in sediments aged between 2.1 and 0.8 Ma, mainly within interglacials.
The strong likelihood that the rise sediments had recorded a history of periodic ice sheet advance to the continental shelf edge during glacial cycles, throughout the period examined, has prompted a wide range of detailed postcruise measurements and analyses of sediment features, including IRD (Cowan, Chap. 10, and Hassler and Cowan, Chap. 11, both this volume), grain size (Pudsey, Chap. 12, and Moerz and Wolf-Welling, Chap. 24, both this volume), spectral reflectance (Wolf-Welling et al., Chap. 21, this volume), and a range of other sedimentological, geochemical, and mineralogical parameters (Hillenbrand and Ehrmann, Chap. 8; Hillenbrand and Fütterer, Chap. 23; Kyte and Vakulenko, Chap. 4; Pudsey, Chap. 25; Wolf-Welling et al., Chap. 15, all this volume). Some of these contributions include interpretations of aspects of regional climatic or sedimentologic evolution, but by no means all; their data are in part a substantial resource on which further interpretive work can be based. Hardly any have had access to the consensus stratigraphy or the revised mcd scales noted above.
Two studies have undertaken spectral analyses of downhole log or core parameters. Lauer-Leredde et al. (Chap. 32, this volume) have analyzed several parameters from downhole logs (thorium/potassium ratio and natural gamma at Site 1095 and uranium, neutron-derived porosity, and natural gamma at 1096) and cores (magnetic susceptibility and bioturbated intervals as described on board ship at Site 1095 and chromaticity parameter L* at Site 1096) over much or all of the depths of the deepest hole at each site, and Pudsey (Chap. 25, this volume) has included analysis of core magnetic susceptibility and chromaticity parameter a* over three short intervals (0-0.4, 2-3, and 3.7-4.3 Ma) at Site 1095. The results are ambiguous. A wide range of spectral peaks of similar amplitude was found, some of which were interpreted as caused by orbitally induced ("Milankovich") insolation changes, others not. The authors have ignored likely diagenetic effects and have recognized that imperfections in the shipboard stratigraphy could affect the analysis. However, for whatever reason, the orbital cyclicity did not stand out; the questions of when and how the Antarctic Peninsula (or Antarctic) ice sheet and adjacent Southern Ocean responded to orbital forcing are extremely important and remain open (and are discussed further below).
Hillenbrand and Ehrmann (Chap. 8, this volume) have used clay mineralogy at Sites 1095 and 1096 to examine the long-term variation in ice sheet behavior. Using a model of the modern glacial-interglacial variation in sediment provenance initiated on board ship (Barker, Camerlenghi, Acton, et al., 1999) and supported by piston core studies by Pudsey (2000) and Hillenbrand (2000) in which mainly downslope, direct transport of sediment high in chlorite during glacials alternates with mainly along-slope (southwestward), indirect transport of sediment high in smectite during interglacials (e.g., Fig. F7), they conclude that periodic migration of the Antarctic Peninsula ice sheet to the continental shelf edge has persisted for the past 9 m.y. with little change. Pudsey (Chap. 25, this volume) reports the presence of poorly sorted sand (most probably IRD), mainly within the bioturbated facies, throughout the section at Site 1095, in agreement with these results. Cowan (Chap. 10, this volume) describes time variations in IRD (sand fraction) abundance over the past 3.1 m.y. at Site 1101, which are difficult to interpret because non-IRD accumulation rates may have varied. Cowan sees cyclic IRD abundance, with peaks in the later glacial and succeeding interglacial parts of cycles (based on the shipboard lithologic definition of glacial cyclicity), and particularly prominent peaks at 0.88, 1.9, and 2.8 Ma. She identifies those cycles with orbital cyclicity from ~1.9 Ma on. Hassler and Cowan (Chap. 11, this volume) investigate IRD (pebble) shape and provenance over the same period, using samples from all drift sites in order to increase total sample number. They suggest a change from mainly englacial and supraglacial IRD transport to ice-base transport at ~0.76 Ma. So late a change in implied glacial state is at odds with most other data and interpretations from the leg.
Hillenbrand and Fütterer (Chap. 23, this volume) have examined biogenic opal (mainly from diatoms) in samples from all of the rise drift sites as a measure of paleoproductivity and, by implication, of ocean temperature within the photic zone. They infer a moderately warm late Miocene, a warm early Pliocene, and a cooling between ~3.1 and 1.8 Ma (late Pliocene) to present-day conditions (Fig. F8) and see a sedimentation rate influence on biogenic silica preservation. They assume that sea-ice cover has been a controlling factor throughout, in contrast to Pudsey (Chap. 25, this volume), who concludes (from a less precise but similar diatom count at Site 1095) that sea ice was unimportant before the Pleistocene because diatoms occurred in both glacial and interglacial facies at the site. Essentially, both studies show the same evolution of shallow ocean temperature at this margin, which is largely compatible also with studies of Southern Ocean climate change (e.g., Abelmann et al., 1990; Kennett and Barker, 1990; Hodell and Venz, 1992; Bohaty and Harwood, 1998) and generally corresponds to global climate change. However, the question of sea-ice cover is one aspect of a concern crucial to Antarctic paleoclimate; in examining features of drift deposition, it is necessary to distinguish between an oceanic and an ice sheet response to orbital forcing. This is discussed further below.
The carbonate preserved at Sites 1096 and 1101 between 2.1 and 0.8 Ma was not interpreted on board ship as an indication that the Polar Front had migrated south of the sites, or disappeared, but its modern prominence in sediments north of this feature was noted. Winter and Wise (Chap. 26, this volume) caution against inferences of Polar Front migration or of specific paleotemperatures from nannofossil preservation. Oxygen and carbon isotopic ratios have been measured on Neogloboquadrina pachyderma (sinistral), by far the most common foraminifer (Barker et al., Chap. 20, this volume), demonstrating the potential of the method in this environment (and, incidentally, showing little or no evidence of elevated shallow-water temperatures). Although most abundant during interglacials, fresh foraminifers (almost always monospecific) are present in sufficient numbers throughout several glacial cycles to allow a single cycle to be examined in detail. For the period between 0.7 and 2.1 Ma, Guyodo et al. (2001) show a convincing correlation of magnetic susceptibility at Site 1101 and the oxygen isotopic record of orbitally induced climate change from ODP Site 677, supporting the view that the interhemispheric responses to orbital insolation change were at that time approximately in phase (a strong correlation between magnetic susceptibility, chromaticity parameter b*, and carbonate content over that period at Site 1101 had been identified on board ship; see fig. F8 of Shipboard Scientific Party, 1999d).
The unusual porosity-depth characteristics at the rise drift sites had been noted on board ship (Barker, Camerlenghi, Acton, et al., 1999), although physical properties data from different sources did not always agree. Volpi et al. (Chap. 17, this volume) have reanalyzed the available porosity, bulk density, and seismic velocity information. They identify a combination of measurements on individual samples, vertical seismic profiles, and acoustic tomography data as the optimal data set for use in calibrating the seismic reflection data set from the continental rise and confirm in passing the anomalous (underconsolidated) nature of the drift sediments.
The possibility that the Eltanin asteroid impact (Kyte et al., 1981, 1988), which occurred at ~2.15 Ma (Pliocene) some 1300 km away in the southeast Pacific, affected sedimentation on the rise drifts (Barker and Camerlenghi, 1999) has been pursued by a search for high Iridium concentrations in selected samples from Holes 1096B and 1096C (Kyte, Chap. 9, this volume). No strong evidence has been found, but the magnetobiostratigraphic synthesis (Iwai et al., Chap. 36, this volume) verifies that the sampled intervals at Site 1096 were correctly identified.